The impact of natural versus anthropogenic aerosolson atmospheric circulation in the CommunityAtmosphere Model

Robert J. Allen • Steven C. Sherwood

Received: 14 October 2009 / Accepted: 17 August 2010

� The Author(s) 2010. This article is published with open access at Springerlink.com

Abstract The equilibrium response of atmospheric cir-

culation to the direct radiative effects of natural or

anthropogenic aerosols is investigated using the Commu-

nity Atmosphere Model (CAM3) coupled to two different

ocean boundary conditions: prescribed climatological sea

surface temperatures (SSTs) and a slab ocean model.

Anthropogenic and natural aerosols significantly affect the

circulation but in nearly opposite ways, because anthro-

pogenic aerosols tend to have a net local warming effect

and natural aerosols a net cooling. Aerosol forcings shift

the Intertropical Convergence Zone and alter the strength

of the Hadley circulation as found in previous studies, but

also affect the Hadley cell width. These effects are due to

meridional gradients in warming caused by heterogeneous

net heating, and are stronger with interactive SST. Aerosols

also drive model responses at high latitudes, including

polar near-surface warming by anthropogenic aerosols in

summer and an Arctic Oscillation (AO)-type responses in

winter: anthropogenic aerosols strengthen wintertime zonal

wind near 60�N, weaken it near 30�N, warm the tropo-

sphere, cool the stratosphere, and reduce Arctic surface

pressure, while natural aerosols produce nearly opposite

changes. These responses are shown to be due to modu-

lation of stratospheric wave-driving consistent with

meridional forcing gradients in midlatitudes. They are

more pronounced when SST is fixed, apparently because

the contrast in land-ocean heating drives a predominantly

wavenumber-2 response in the northern hemisphere which

is more efficient in reaching the stratosphere, showing that

zonal heating variations also affect this particular response.

The results suggest that recent shifts from reflecting to

absorbing aerosol types probably contributed to the

observed decadal variations in tropical width and AO,

although studies with more realistic temporal variations in

forcing would be needed to quantify this contribution.

Keywords Arctic Oscillation � Aerosol � Hadley cell �Atmospheric circulation

1 Introduction

Aerosols are important to the climate system because they

affect the radiative balance of the planet. First, they scatter

and absorb solar radiation, and therefore contribute to

atmospheric solar heating and surface cooling (Ramana-

than et al. 2001a). This is referred to as the aerosol direct

effect. Because absorbing aerosols heat the atmosphere,

they can also decrease relative humidity and evaporate

clouds (Ackerman et al. 2000). This is referred to as the

semi-direct effect and can cause additional heating of the

surface. Aerosols also have an impact on cloud micro-

physical properties (and hence, the radiative properties).

Because aerosols act as cloud condensation nuclei (CCN),

an increase in aerosols leads to an increase in CCN, which

in turn leads to an increase in cloud droplet number den-

sity, a reduction in cloud droplet effective radius and

ultimately, brighter clouds (Twomey 1977). This is refer-

red to as the first aerosol indirect effect, which enhances

reflection of solar radiation back to space by clouds (and

also reduces the solar radiation reaching the ground). The

smaller cloud droplets also promote suppression of

R. J. Allen (&)

Department of Earth System Science,

University of California Irvine, Irvine, CA 92697, USA

e-mail: rjallen@uci.edu

S. C. Sherwood

Climate Change Research Centre,

University of New South Wales, Sydney, Australia

123

Clim Dyn

DOI 10.1007/s00382-010-0898-8

precipitation. This is referred to as the second indirect

effect and leads to an increase in cloud lifetime (Albrecht

1989).

Several studies have documented the climatic impacts of

reflecting [e.g. sulfate and organic carbon (OC)] aerosols.

For example, Roeckner et al. (1999) use a coupled general

circulation model (GCM) with a tropospheric sulfur cycle

and find that increasing anthropogenic sulfate aerosols

mitigate greenhouse warming by reflecting shortwave

radiation. They also find the intensity of the global

hydrological cycle becomes weaker if direct and indirect

sulfate aerosol effects are included. Moreover, the simu-

lated temperature trends during the last half of the twen-

tieth century are in better agreement with observations if

both greenhouse gases and sulfate aerosols are included

(e.g. Mitchell et al. 1995).

Sulfate aerosols can affect the large-scale atmospheric

circulation, particularly in the tropics, causing a southward

shift of the intertropical Convergence Zone (ITCZ) (Wil-

liams et al. 2001). This shift is attributed to a reduced inter-

hemispheric temperature gradient since reflecting aerosols

predominately cool the NH. Rotstayn and Lohmann (2002)

suggests that this shift may have contributed to the Sahe-

lian drought of the 1970s and 1980s. Similarly, Cox et al.

(2008) finds that sulfate aerosols alter the meridional sea

surface temperatures (SST) gradient in the equatorial

Atlantic, which affects Amazonian precipitation; future

aerosol reductions imply increasing risk of Amazonian

drought.

Reflecting aerosols are also capable of affecting high-

latitude circulation. Several studies (e.g. Robock 2000; e.g.

Shindell et al. 2004) have shown that tropical volcanic

aerosols (which are primarily composed of sulfate) affect

NH high-latitude winter climate, causing a positive Arctic

Oscillation (AO) anomaly. The conventional mechanism

explaining this effect is heating of the low latitude strato-

sphere (and ozone loss at high latitudes, which cools the

polar lower stratosphere) enhances the meridional tem-

perature gradient and strengthens the westerly winds near

the tropopause. The enhanced westerlies then propagate

down to the surface via wave-mean flow interactions,

creating a surface temperature response pattern typical of

the AO. Stenchikov et al. (2002) also finds a positive AO

response results from low latitude cooling of the tropo-

sphere (since volcanic aerosols predominantly reflect solar

radiation).

Recent research has focused on the impact of absorbing

aerosols on regional climate, especially in the South Asian

region, where emissions of fossil fuel black carbon (BC)

increased by *6-fold since 1930 (Ramanathan et al.

2005). Based on equilibrium model experiments with

prescribed SSTs (as well as prescribed aerosol forcing),

Chung et al. (2002) and Menon et al. (2002) show that

absorbing aerosols lead to enhanced low-level convergence

and precipitation over the haze area, including India and

China. Bollasina et al. (2008) reach a similar conclusion,

finding that in years when more aerosols are observed in

May, cloud and precipitation amounts are reduced,

increasing the surface temperature through solar heating.

This results in a stronger June/July monsoon. Similarly,

Randles and Ramaswamy (2008) find that at higher

extinction-optical-depths, enhanced low-level convergence

and ascent overcome the stabilizing effects of absorbing

aerosols and strengthens the monsoon. However, pre-

scribed increases in scattering aerosols alone has the

opposite effect. A weakening of the monsoon is obtained

either with prescribed SSTs (Chung and Ramanathan 2006;

Ramanathan et al. 2007) or when prognostic SSTs in a

coupled model can respond to the aerosol forcing (Rama-

nathan et al. 2005; Meehl et al. 2008). This weakening is

caused by reduced evaporation over the Indian Ocean, a

reduced land–sea temperature contrast and meridional SST

gradient in the northern Indian Ocean, and increased

atmospheric stability.

A few studies have identified broader impacts of

absorbing aerosol forcing. Wang (2004) uses an interactive

aerosol-climate model based on equilibrium simulations

with CCM3 and a slab ocean model to examine the direct

radiative forcing of BC aerosols. He finds that BC causes a

significant change in tropical convective precipitation, with

a strengthened (weakened) Hadley cell in the Northern

(Southern) Hemisphere. These results are similar to Rob-

erts and Jones (2004) and Yoshimori and Broccoli (2008).

Wang (2004), however, concludes that the effects of BC

aerosols on climate are more significant at the regional,

rather than the global, scale. Because he assumes BC is

externally mixed (physically separate from other particles)

the aerosol effects on climate are likely underestimated

(Ramanathan and Carmichael 2008).

To address the uncertainty in the mixing state of BC,

Chung and Seinfeld (2005) conduct separate GCM experi-

ments assuming BC is both externally and internally

mixed. They find significantly more warming with inter-

nally mixed BC, 0.54 versus 0.29 K for the NH annual

mean surface air temperature. Similarly, internally mixed

BC yields enhanced atmospheric circulation changes, with

a larger northward shift of the ITCZ, and larger increases

(decreases) in precipitation between 0 and 20�N (0–20�S).

Relatively little attention has been devoted to global-

scale/remote responses to absorbing aerosols. Kim et al.

(2006) find that absorbing aerosols (dust and black carbon)

excite a planetary-scale teleconnection pattern in sea level

pressure, temperature, and geopotential height spanning

north Africa through Eurasia to the north Pacific. They also

show that the surface temperature signature associated with

the aerosol-induced teleconnection is very similar to the

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

spatial pattern of observed long-term trend in surface

temperature over Eurasia. Similarly, Chung and Rama-

nathan (2003) investigate the influence of the interannual

variability of the SE Asian haze on global climate. Two

CCM3 experiments are conducted, representing extreme

locations of the forcing. The remote impacts of the two

experiments are significantly different, with the extended

haze experiment yielding zonal wind changes similar to

those associated with the positive AO (the shrunk haze

experiment yields negligible changes). The SE Asian haze

may therefore partially explain the the observed increase in

AO variability (Feldstein 2002). In addition to interannual

variability, aerosols have also exhibited relatively large

trends over the latter half of the twentieth century, with

increasing global emissions of BC (although they have

fallen somewhat since 1990) (Novakov et al. 2003; Ito and

Penner 2005; Bond et al. 2007) and decreasing emissions

of sulfate aerosols since the 1970s (van Aardenne et al.

2001; Smith et al. 2004). A recent study attributes up to

75% of the Arctic surface temperature increase since 1976

to aerosols, specifically the decrease in sulfate in North

America and Europe and increase in BC from SE Asia

(Shindell and Faluvegi 2009). Relatively little is known

about how such aerosol trends have affected atmospheric

circulation.

The above studies compose a complicated and uncertain

overall picture of the regional and global climate impacts

of aerosols, based on a variety of aerosol representations

and boundary conditions. Some uncertainty stems from a

lack of knowledge about the spatial and temporal evolution

of aerosol emissions, especially BC and organic carbon

(Bond et al. 2007). Additional uncertainty comes from

issues in how to model the aerosol effect. In most GCMs,

the absorbing aerosol forcing (primarily BC) is underesti-

mated by a factor of 2–4 (Sato et al. 2003; Koch et al.

2009). This underestimation is due to a number of factors,

including neglect of the internally mixed state of BC with

other aerosols, neglect of biomass burning (about 40% of

total BC emission) and BC concentrations that peak too

close to the surface (Ramanathan and Carmichael 2008).

There is also uncertainty in the the relative amounts of

different aerosol types. Finally, studies make different

assumptions as to the behavior of SST, which is clearly

important in that cooling the surface (at least in the Tro-

pics) is a key avenue by which aerosols can alter the

circulation.

In this study, we simulate the direct effect of aerosols

using the Community Atmosphere Model version 3 (CAM3,

Collins et al. 2004). We bring some clarity to the role of

direct aerosol forcing in driving circulation changes by using

recent observationally derived forcing estimates, which are

significantly higher than those used in most previous model

studies; by examining the sensitivity to ocean behavior and

aerosol properties, which can significantly affect atmo-

spheric responses; and by analyzing aspects of the global

circulation response that have not been widely examined. For

simplicity, we examine only equilibrium responses, and

ignore microphysical (indirect) effects on clouds. The paper

is organized as follows: Sect. 2 discusses the aerosol forcing

estimate used in this analysis and the experiment design for

the global modeling. Section 3 presents the results, including

circulation changes at low latitudes (e.g. mean meridional

mass circulation changes) and mid/high-latitudes (zonal

wind changes and the AO). Conclusions follow in Sect. 4. A

companion paper (Allen and Sherwood 2010) focused on the

effect of aerosol heating on cloud distributions and conse-

quent ‘‘semi-direct’’ effects on global climate.

2 Experimental design

2.1 Satellite derived aerosol forcing

We use the satellite-based aerosol forcing of Chung et al.

(2005) (hereafter, CH05), which is based on the Moderate

Resolution Imaging Spectroradiometer (MODIS), with gaps

and errors filled and corrected using the Georgia Tech/

Goddard Global Ozone Chemistry Aerosol Radiation and

Transport (GOCART) model, as well as ground based

observations from the AErosol RObotic NETwork (AER-

ONET). The anthropogenic fraction of the aerosol forcing is

obtained by differencing the natural and total aerosol

forcing. The natural forcing is obtained from two inde-

pendent sources. The standard estimate (and the one we use)

is simply the prediction of the GOCART model (Chin et al.

2002). The second approach—which provides a lower

bound on the natural fraction—uses the ratio of MODIS

aerosol optical depth (AOD) from larger aerosol particles to

that of all particles (since anthropogenic aerosols are mostly

sub-micron sized). The single scattering albedo (SSA) and

asymmetry parameter are also derived from GOCART and

adjusted based on AERONET observations. These param-

eters are then introduced into the Monte-Carlo Aerosol

Cloud Radiation (MACR) model (Podgorny et al. 2000)

with ISCCP (International Cloud Climatology Project) D2

cloud data to derive, for the years 2000–2003, the aerosol-

induced reduction of solar radiation at the surface (FSFC)

and increase in atmospheric solar absorption (FATM). We

note that the aerosol forcing was likely different (i.e., larger

SSA due to a larger ratio of sulfate to carbon) in earlier

times (Smith et al. 2004; Bond et al. 2007).

The global, annual mean anthropogenic aerosol forcing

thus obtained is 3.1 W m-2 in the atmosphere (SSA ranges

from 0.85 to 0.88) and -3.5 W m-2 at the surface. Based on

sensitivity experiments CH05 estimated the corresponding

uncertainties at 20 and 10%, respectively. For a given

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

aerosol type, the major sources of uncertainty are the SSA

and the vertical distribution of the aerosols. For example, if

all aerosols are confined below 1 km (2–4 km), atmo-

spheric absorption decreases (increases) to 2.7 (3.2) W

m-2. There is also uncertainty in distinguishing natural

from anthropogenic aerosols. If the MODIS fine mode

fraction (ratio of small mode to total optical depth) is used

for the anthropogenic fraction over the ocean (instead of

GOCART estimates), atmospheric forcing decreases to

3.0 W m-2. There may also be a bias, however, in MODIS

retrievals of this ratio, which have been found to be sys-

tematically larger than in situ measurements by 0.2 in the

Aerosol Characterization Experiment (ACE)-Asia region

(Anderson et al. 2005). This underestimation of retrieved

particle size distributions is consistent with an analysis of

the Puerto Rico Dust Experiment (PRIDE) data set (Levy

et al. 2003). These results suggest a possible overestimation

of satellite-derived anthropogenic aerosol forcing.

The forcing estimates used here are similar to others

recently developed from satellite and other observations. In

terms of the top of the atmosphere (TOA) direct radiative

forcing under cloudy skies, Yu et al. (2006) obtains

-0.5 ± 0.33 W m-2, compared to CH05’s -0.35 ± 0.25.

In terms of all-sky conditions (i.e., no clouds) Bellouin

et al. (2005) obtains -0.8 ± 0.2 W m-2 compared to

CH05’s -1.08 W m-2. For atmospheric forcing under

clear skies, Bellouin et al. (2005) obtains 2.5 W m-2 (all

aerosols), compared to CH05’s estimate of 3.4 W m-2

(anthropogenic aerosol only). Thus, aerosol forcings now

seem to be known to an accuracy of 50% or better globally.

Importantly, these estimates are several times larger than

those derived previously from models, which average 1.24 W

m-2 for clear-sky conditions (Forster et al. 2007), indicating

that many published model results may significantly under-

estimate some aspects of aerosol-forced responses.

2.2 Model experiment details

It is not practical to run idealized/equilibrium aerosol

forcing simulations with full ocean models due to the great

expense of reaching equilibrium in the ocean with each

aerosol scenario. The remaining options are to prescribe

SSTs (thus preventing any ocean-modulated response), or

use a slab ocean model (which often exaggerates local

responses to surface flux anomalies because ocean heat

transports cannot change). To assess the robustness of our

results to ocean behavior, we repeat otherwise identical

experiments with both of these: a ‘‘data ocean model’’

(DOM) (i.e., fixed climatological seasonal cycle of SSTs)

and a slab ocean-thermodynamic sea ice model (SOM). To

our knowledge this has been done before only by Wang

(2004), who did not show any of the DOM results. With the

SOM, the ocean mixed layer contains an internal heat

source (Q flux) representing seasonal deep water exchange

and horizontal ocean heat transport. SOM (DOM) experi-

ments are run for 150 (100) years, the last 100 (70) of

which are used in this analysis, during which there was no

significant trend in TOA net energy flux. The DOM and

SOM setups constitute oceans with infinite and (at equi-

librium) zero local heat capacity, respectively, which we

argue is likely to bracket the behavior that would occur

with a more realistic ocean. While anomalous ocean heat

transports could cause SSTs in some regions to change less

or even in the opposite direction to what a slab model

would predict, the predominant response is likely to be at

least qualitatively captured in most places, as suggested for

example by modeled responses to aerosol forcing (Evan

et al. 2009). The DOM simulation, on the other hand, may

be representative of the response to forcing variations on

intraseasonal and shorter time scales where the ocean

response would be muted by thermal inertia.

Monthly aerosol forcings (FATM and FSFC)—as esti-

mated by CH05—are incorporated into the CAM radiation

module. Aerosol absorption is uniformly distributed

throughout the lowest *3 km of the atmosphere, in accord

with observation (Ramanathan et al. 2001b). Although

aerosol forcing is almost independent of zenith angle (h)

when h is small, aerosol forcing! 0 as h! 90�. Thus, the

added aerosol forcing is multiplied by a scaling factor that

depends on h (Chung 2006). Although prescribed aerosol

forcing is not allowed to change in response to meteoro-

logical (e.g., wind, precipitation) changes, Chung (2006)

shows that this approach yields similar results (i.e., changes

in precipitation) as the on-line approach, where the distri-

bution and amount of aerosols in the atmosphere are pre-

dicted based on emissions (e.g., Wang 2007). CAM is run

at T42 resolution (*2.8� 9 2.8�) with 26 vertical levels.

Table 1 lists the CAM experiments performed, as well

as the five signals (i.e., differences between experiments)

that will be discussed. The anthropogenic ? natural

Table 1 Definition of the three CAM aerosol experiments and

signals (differences between experiments)

Experiment Aerosols used

NULL None

NATURAL Natural

ALL Anthropogenic ? natural

Signal Experiments used

NAT NATURAL–NULL

ALL ALL–NULL

ANTHRO ALL–NATURAL

Each experiment was run with a DOM (fixed SST) and SOM (slab

ocean)

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

experiment is compared to two control runs, the natural

experiment (to get the anthropogenic signal, designated as

ANTHRO) and a control run with no aerosols (to get the

total aerosol signal, designated as ALL). The natural aero-

sol signal (designated as NAT) is obtained by comparing

the natural experiment with the no aerosol control run.

CAM default aerosol are removed by setting the aerosol

mass mixing ratio for the five aerosol species (carbona-

ceous, dust, sea salt, sulfur, and volcanic) to zero.

To help interpret the results (in particular the importance

of the aerosol spatial heterogeneity), two additional CAM

SOM experiment are performed. Both use default CAM

settings (e.g. aerosols), except a globally uniform heating

source of 0.1 K day-1 (about 3.5 W m-2, similar to FATM)

is added to the lowest *3 km of the atmosphere (as with

ANTHRO) in the experiment; the control is analogous, but

has no heating source. The difference between these two

experiments (i.e. the signal) is designated as UHT (uniform

heating). Results are based on the last 30 years of a 70 year

integration.

Figure 1 shows the prescribed zonal mean DJF and JJA

atmospheric heating (FATM) and reduction in surface solar

radiation (FSFC) for the three experiments. The forcing is

largest in the NH, where the major sources of aerosols

exist. There is also a clear seasonal signal, with larger

forcing during each hemisphere’s summer, when solar

radiation is maximum. ANTHRO FATM is much larger than

that for NAT, due to a greater proportion of absorbing

aerosols. Based on anthropogenic aerosols only, the SSA is

0.85–0.88; adding natural aerosols significantly decreases

the absorption, increasing the total SSA to 0.93–0.95

(Chung et al. 2005).

During DJF, the maximum forcing occurs near 10�N,

primarily due to anthropogenic aerosols from central

Africa (due to biomass burning) and India (due to fossil-

fuel burning) and natural aerosols (dust) from central

Africa. ANTHRO FATM (FSFC) peaks near 6(-7) W m-2.

The NAT forcing is substantially less, with FATM (FSFC)

peaking near 0.5(-3) W m-2. Because ANTHRO FATM is

nearly equal in magnitude to FSFC, ANTHRO is comprised

of a high proportion of absorbing aerosols. However,

because NAT FATM is only *20% of FSFC, a low pro-

portion of absorbing aerosols (i.e., a high proportion of

reflecting aerosols) comprise NAT.

During JJA, ANTHRO FATM for the entire NH is at least

3 W m-2. Two distinct maxima exist at 5�S and 40�N, with

FATM approaching 7 and 8.5 W m-2, respectively. These

maxima are due to aerosols from southern Africa and

southeast Asia/China. NAT has substantially less heating,

with most of the NH FATM between 0.5 and 1.0 W m-2,

peaking near 15�N at 1.2 W m-2. India, the Middle East,

the Gobi Desert, and parts of China contribute most to

NAT, again primarily due to dust. In terms of FSFC, the

disparity between anthropogenic and natural aerosol is

much less. ANTHRO (NAT) minimizes at *-9(-5) W

m-2 near 40(15)�N. Similar to DJF, ANTHRO (NAT) JJA

Fig. 1 DJF (top) and JJA

(bottom) zonal mean reduction

in surface solar radiation (FSFC;

left) and atmospheric heating

(FATM; right) for each

experiment. Units are W m-2

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

possess a much higher (lower) proportion of absorbing

versus reflecting aerosols. This is true year-round.

Note that the atmospheric absorption by anthropogenic

aerosols approximately balances the surface cooling—that

is, the net effect is to move energy from the ocean to the

atmospheric boundary layer with little net power input.

This will heat the system overall when SSTs are fixed, but

with a slab ocean, net fluxes into the surface cannot change

over time so the net system heating will be small. By

contrast, natural aerosols cool the surface without heating

the atmosphere much. This will produce a net power sink

over land areas and oceans represented by a slab model, but

will have no effect over the DOM ocean. Thus the overall

power input by aerosols and its land–ocean contrast will

depend on both the aerosol type and the assumed ocean

behavior.

3 Results

3.1 Vertical temperature structure

Figure 2 shows the vertical cross section of temperature

change (DT) in DJF and JJA, for both ocean conditions and

aerosol types. In most of the experiments the troposphere

warms, although natural aerosols (NAT) produce little

warming with fixed SST (DOM) and significant cooling

with the slab ocean (SOM). This accords with the differ-

ences in net forcing between SOM and DOM, with the

surface energy sink well exceeding the atmospheric source

in the DOM case, cooling the oceans and thus the tropo-

sphere. Low latitudes tend to warm the most in DJF,

broadly consistent with the location of maximum aerosol

heating at that time (Fig. 1), although ANTHRO warming

peaks over Antarctica and near 40�N/S. Because the peak

heating is near 10N, the midlatitude peaks must be driven

by changes in the circulation (see Sect. 3.3). The Antarctic

warming is probably driven by local forcing due to strong

summertime insolation. In JJA the tropospheric changes

are strongest in the northern hemisphere where the forcing

is largest; in particular there is strong warming in JJA in the

Arctic lower troposphere.

In the NAT experiments, the stratosphere generally

warms despite the lack of any direct forcing there. The

reason for this is unclear. No similarly broad change occurs

with ANTHRO forcing, but strong localized changes occur

in the polar regions where the stratospheres tend to cool

with ANTHRO forcing but warm with NAT. These changes

are stronger in the Arctic than the Antarctic, and stronger in

local winter than in summer. Since there is no direct forcing

in the stratosphere, these changes must be associated with

stratospheric dynamical responses (Sect. 3.3).

It is interesting to compare more carefully the slab ocean

and fixed SST experiments. They produce similar DT in

most places despite the different ocean response to surface

forcing. But a significant exception is that the wintertime

polar stratospheric (and to a lesser extent tropospheric)

changes noted above are much stronger in the DOM (fixed

SST) runs than in the SOM runs. This difference is much

larger in the Arctic winter than the Antarctic one, and is

greater for ANTHRO forcing than for NAT. The larger

difference in DOM suggests that land-ocean contrasts may

play a role in driving the dynamical responses (Sect. 4).

Another somewhat surprising result is that, with ANTHRO

forcing in DJF, the tropical troposphere warms at least as

much in SOM as in DOM; in the other three cases, DOM

warms more, as would be expected from the lack of ocean

cooling. Moreover, even in JJA the slab ocean has signif-

icantly less impact on tropical warming for ANTHRO than

for NAT, despite the surface heat sink being similar for

both aerosol forcings. The reasons for this are not clear but

may again have to do with differing dynamical responses

with the two oceans, as shown by evidence noted above.

The large Arctic near-surface warming is interesting in

light of its possible relevance to recent trends (c.f. Shindell

and Faluvegi 2009). In these experiments it is enhanced by

snow and ice melt (not shown) and moreover by decreased

low cloud amounts over land areas (Allen and Sherwood

2010), both of which lead to increased solar heating.

Despite an applied surface forcing of only -2 W m-2

(Fig. 1) at high NH latitudes, the equilibrium change in net

short wave radiation at the surface is nearly 10 W m-2 for

both ANTHRO runs. It is interesting that this warming is

not significantly reduced with a slab ocean, which is

probably because of the large fraction of land at high

northern latitudes plus the fact that ocean circulation

changes cannot occur in either experiment.

Compared to other absorbing aerosol studies (Chung

and Seinfeld 2005; Yoshimori and Broccoli 2008), the

vertical structure of warming for ANTHRO SOM is simi-

lar, but does not extend as far vertically. Chung and

Seinfeld (2005) find maximum BC-induced warming

(internally mixed aerosols) near 400 hPa at 30�N, espe-

cially during JJA, of about 1.2 K . The corresponding

warming in our study is *0.4 K. The lesser warming

found here is likely because ANTHRO SOM is forced with

not only absorbing (i.e., BC) aerosols, but other anthro-

pogenic aerosols, like sulfate, that tend to cool the surface

and atmosphere. The global (NH) annual mean DTsfc is

0.09(0.18) K, compared to Chung and Seinfeld’s DT sfc of

0.37(0.54) K. However, global cooling in ANTHRO SOM

DT sfc is similar to that of Wang (2004), who assumes

externally mixed BC aerosols, at 0.09 K for the global

annual mean.

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

3.2 Hadley cell/ITCZ

Figure 3 shows the change in the mean meridional mass

circulation (DMMC), defined by calculating the northward

mass flux above a particular pressure level (i.e., a mass

streamfunction). The clearest changes are to the Hadley cell,

which is not surprising given that the forcings peak near the

latitude of the ITCZ (Fig. 1). In both seasons and for both

forcings, these are larger with the slab ocean. Furthermore,

the changes are opposite for ANTHRO and NAT.

In DJF with a slab ocean, ANTHRO forcing weakens

the MMC, while NAT strengthens it. The ANTHRO

Fig. 2 DJF (top four panels)

and JJA (bottom four panels)

temperature change. Symbolsrepresent significance (assessed

with a standard t test using the

pooled variance) at the 90%

(diamond), 95% (cross) and

99% (dot) confidence level.

Units are K

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

weakening also appears in DOM but not strongly. Opposite

changes occur in JJA, with only the NAT signal appearing

(weakly) in DOM. For ANTHRO forcing, this strength-

ening/weakening is associated with a northward shift of the

tropical MMC maxima; NAT forcing shifts the MMC

maxima southward. Table 2 shows that the MMC maxima

shifts *0.5–1.0� northward (southward) for ANTHRO

(NAT) SOM. These results are consistent with other slab

Fig. 3 DJF (top four panels)

and JJA (bottom four panels)

mean meridional mass

circulation change (colorshading). Negative (positive)

values represent

counterclockwise (clockwise)

circulation change. Symbolsrepresent significance as in

Fig. 2. Units are 1010 kg s-1.

Also included is the

climatological MMC (contour

lines) based on the

corresponding control

experiment. Contour interval is

5 9 1010 kg s-1, with negative

values dashed

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

ocean aerosol studies focusing on the direct effects of BC

aerosols (Wang 2004; Roberts and Jones 2004; Chung and

Seinfeld 2005; Wang 2007; Yoshimori and Broccoli 2008),

and the direct (Yoshimori and Broccoli 2008) and indirect

(Rotstayn et al. 2000; Williams et al. 2001) effects of

sulfate aerosols.

The change in MMC is associated with the change in

inter-hemispheric SST difference, such that anomalous

heat is transported from the warmed to the cooled hemi-

sphere (Zhang and Delworth 2005; Broccoli et al. 2006;

Mantsis and Clement 2009). Because most of the aerosol

forcing is north of the equator, the associated DT is larger

there too. ANTHRO aerosols warm the NH troposphere

(Fig. 2), thus weakening the inter-hemispheric temperature

difference in DJF and strengthening it in JJA. The resulting

annual mean DSST is 0.22 K in the northern hemisphere

and 0.02 K for the SH, in the slab ocean simulations.

Because NAT aerosols cool the NH troposphere (Fig. 2),

their effects are generally opposite to those of ANTHRO.

For NAT SOM, the NH annual mean DSST is -0.28 K,

compared to -0.12 K in the SH. This argument is also

consistent with the response (not shown) of the uniform

heating experiment (UHT), where the MMC weakens in

both DJF and JJA. This is because the SST warming is

greater in the winter hemisphere in both seasons, leading to

a weaker inter-hemispheric SST gradient. This also leads to

northward displacement of the tropical MMC maxima

(Table 2) during DJF (like ANTHRO), but southward

displacement during JJA (like NAT).

In our experiments, the strength of the inter-hemispheric

SST and the cross-equatorial heat transport (calculated as

the product of the meridional velocity and temperature, VT)

are closely related. The correlation of the interhemispheric

DSST and zonal (equatorial) mean DVT (averaged from

850 to 150 hPa) over three seasons (JJA, DJF and ANN)

and four signals (NAT and ANTHRO for DOM and SOM)

is -0.93. Similarly, the corresponding correlation between

the zonal (equatorial) mean DVT and DMMC (evaluated at

500 hPa) is 0.99. Thus, an increase in interhemispheric SST

decreases the northward heat transport by modifying the

MMC. Similar results are obtained with different defini-

tions of the MMC (e.g., using other pressures or tropical

latitudes) or by looking at DT700 rather than DSST. The

changes in interhemispheric DSST, cross-equatorial heat

transport, and MMC change are all smaller in DOM than

SOM simulations.

As would be expected, the above changes are associated

with a shift in the Intertropical Convergence Zone (ITCZ),

reflected in the vertical velocity, convective precipitation,

and cloud fields. The ITCZ moves northward in ANTHRO

and southward in NAT, regardless of season, consistent

with the shift of the tropical MMC maxima. Figure 4 shows

the change in precipitation minus evaporation (P - E) in

the SOM runs for DJF and JJA. The tropical P - E maxima

is displaced northward, relative to the control integration,

for ANTHRO SOM and southward for NAT SOM. Table 2

shows the magnitude of the displacement is up to *1.6�

Table 2 Northward displacement of the tropical (top) MMC and

(bottom) precipitation minus evaporation (P - E) maxima for both

ocean boundary conditions

SOM DOM

DJF JJA ANN DJF JJA ANN

MMC

ANTHRO 0.61** 0.99** 0.18 0.06 0.03 -0.49**

NAT -0.71** -0.49 -0.18 0.0 0.23 0.40

UHT 0.01 -1.2* -0.32 X X X

P - E

ANTHRO 1.6** 0.12 0.79** 0.60 -0.08 0.11

NAT -1.6** -0.16** -0.53** -0.23 0.0 -0.06

UHT 1.4 -0.29* 0.35 X X X

Significance is denoted by bold (C90%); *(C95%) and **(C99%).

Units are degrees latitude. MMC is evaluated at 500 hPa. Precipita-

tion is the sum of both convective and large scale precipitation

Fig. 4 Zonal DJF (top) and JJA (bottom) precipitation minus

evaporation (P - E) change for ANTHRO and NAT SOM. Also

included is P - E from the control experiment. Symbols represent

significance as in Fig. 2. Units are mm day-1

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

during DJF for the SOM experiments, with generally

smaller changes in the DOM runs. The UHT P - E signal is

also consistent with its tropical MMC displacement, with a

northward shift of the tropics during DJF and a southward

shift during JJA.

3.3 Zonal winds

3.3.1 Low-latitude response

Figure 5 shows the change in zonal wind (DU) during the

DJF and JJA seasons. As for the low-latitude MMC

changes, the DU changes are generally opposite for the

ANTHRO and NAT forcings. Unlike DMMC, however,

both surface boundary conditions (DOM and SOM) yield

similar patterns and similar absolute magnitudes of zonal

wind change.

The SOM—and to a lesser extent the DOM—experi-

ments show a meridional displacement of the subtropical

jets. Both the NH and SH jet tends to shift poleward from

ANTHRO forcing, leading to expansion of the Hadley

cell and the tropics; NAT forcing leads to equatorward jet

displacement and contraction of the tropics. The meridi-

onal displacement of the jet (D/jet) was quantified by

interpolating the zonal wind to 0.5� resolution using cubic

splines and then locating the latitude of the maximum NH

and SH zonal wind for each year-month in the upper

troposphere (250–150 hPa). Averaging over season and

taking the difference of the mean jet location (experi-

ment-control) yields the poleward jet displacement, as

shown in Table 3. With both oceans, ANTHRO yields

poleward displacement of the NH jet during every season

except DJF, up to 0.42� during boreal autumn (SON; not

shown). Displacement is smaller and not significant in

DJF as the decrease in wind occurs near the center of the

jet, while the increase occurs nearly 20� poleward of it.

NAT-forced shifts of the NH jet are equatorward, and

again largest for SON at 0.35�. Displacements of the SH

jet are generally in the same direction (poleward for

ANTHRO and equatorward for NAT), but smaller, in

accord with the reduced aerosol forcing. The largest SH

jet displacement occurs for ANTHRO DOM during SON

and for ANTHRO SOM during JJA, where the SH jet

moves poleward by 0.54� and 0.27�, respectively. This

suggests a maximum tropical widening of *0.5� for

ANTHRO SOM (during JJA) and *0.9� for ANTHRO

DOM (during SON); the corresponding maximum con-

traction (NAT JJA) is similar in magnitude at *0.4�. The

seasons of maximum expansion-JJA and SON-corre-

spond to the seasons with maximum aerosol absorption, at

both low latitudes (4.7 and 4.5 W m-2, respectively,

for ±30�), as well as globally. Similar results are obtained

if other upper tropospheric/lower stratospheric pressure

levels are used to measure the jet, or if U is not inter-

polated before locating the jets.

Table 3 shows that alternative measures of tropical

expansion/contraction are consistent with the subtropical

jet displacements. These metrics include (1) the latitude of

the subtropical MMC minima, defined as the NH and SH

latitude where the MMC at 500 hPa becomes zero pole-

ward of the subtropical maxima; and (2) the latitude where

precipitation minus evaporation (P - E) becomes zero on

the poleward side of the subtropical minima (e.g. Hu and

Fu 2007; e.g. Lu et al. 2007; e.g. Johanson and Fu 2009).

Similar to the procedure used to find the jet displacement,

zonal mean MMC and P - E are first interpolated to 0.5�resolution using cubic splines, and then the latitude of

interest is located for each year-month for experiment and

control. Although both metrics yield similar results, we

emphasize the MMC statistic because aerosol semi-direct

effects on clouds, and hence precipitation, likely make the

P - E metric noisier. During DJF, Table 3 shows that

anthropogenic aerosols shift the tropics northward, by

*0.4� for the slab ocean, while natural aerosols shift the

tropics southward by a similar amount. These results are

consistent with the prior section (Table 2), with northward

displacement of the ANTHRO MMC and P - E maxima

(i.e. center of tropics) and southward displacement for

NAT, due to different interhemispheric SST gradient

changes. During JJA, however, both the MMC and P - E

metrics yield tropical expansion for ANTHRO and con-

traction for NAT, of about 0.2� for the slab ocean experi-

ments. Similar tropical expansion for ANTHRO also

occurs during SON. These other metrics are therefore

consistent with the subtropical jet displacement, further

supporting the notion that anthropogenic aerosols promote

tropical expansion, especially during the seasons of larger

aerosol forcing (JJA and SON).

One explanation used to account for tropical expansion/

contraction is that of Held (2000), where the Hadley cell

latitude is proportional to the tropopause height and the

gross dry static stability (potential temperature difference

between the tropopause and surface, hTP - hsfc). Both

Frierson et al. (2007) and Lu et al. (2007) show that a

variety of global warming experiments yield increases in

subtropical static stability, which reduces baroclinic growth

rates and displaces the region of baroclinic instability onset

poleward, resulting in tropical expansion. Similarly, Lor-

enz and DeWeaver (2007) show that tropical expansion is

associated with raising of the tropopause, especially on the

poleward flank of the jet. Lu et al. (2009) attributes the

raising of the tropopause (and consequent tropical expan-

sion) since 1958 to stratospheric cooling caused by

greenhouse gas increases and stratospheric ozone deple-

tion. This is consistent with Haigh et al. (2005), who found

uniform (and high-latitude) heating of the lower

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

stratosphere forces the jets equatorward, while low-latitude

stratospheric heating forces the jets poleward. Equatorward

displacement of the jets in response to a warmer strato-

sphere was also found by Williams (2006).

Our aerosol experiments are generally consistent with

these relationships. For example, Fig. 6 shows that

ANTHRO SOM yields an increase in the JJA tropopause

height/decrease in tropopause pressure [based on a thermal

Fig. 5 DJF (top four panels)

and JJA (bottom four panels)

zonal wind change (colorshading). Symbols represent

significance as in Fig. 2. Units

are m s-1. Also included is the

climatological U (contour lines)

based on the corresponding

control experiment. Contour

interval is 10 m s-1 , with

negative values dashed

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

definition according to Reichler et al. (2003)], whereas

NAT SOM yields a decrease in JJA tropopause height

(other seasons are similar). Furthermore, ANTHRO SOM

yields an increase in JJA gross dry static stability; NAT

SOM yields a decrease. Over all aerosol experiments and

seasons, the correlation between the poleward jet dis-

placement and 30�N/S tropopause height change is 0.76,

significant at the 99% confidence level. Similarly, the

correlation between D/jet and the mid-latitude (30�–50�)

gross dry static stability change is 0.78, also 99% signifi-

cant. The jet displacements are also associated with shifts

of the climatological mid-latitude baroclinicity (dT/dY)

maxima near the jet altitude (300 hPa), with poleward

displacement for ANTHRO and equatorward displacement

for NAT, an expected consequence from thermal wind

balance. However, Fig. 6 shows that these changes tend to

be localized near the latitude of the jets, suggesting a

response, rather than a cause. This is further supported by

comparing DdT=dY to ðD/jetÞ � ðd2T=dY2Þ (not shown),

with the second derivative of temperature based on the

control integration. For UHT, DdT=dY is consistently lar-

ger than ðD/jetÞ � ðd2T=dY2Þ, which suggests the tem-

perature forcing is dragging the jet. In contrast, the aerosol

experiments feature a DdT=dY that is comparable to

ðD/jetÞ � ðd2T=dY2Þ, suggesting displacement of the tem-

perature field along with the jet.

Figure 7, however, shows that D/jet is consistent with

the meridional net aerosol forcing (i.e. FATM ? FSFC)

gradient (dFNET/dY) between the subtropics and midlati-

tudes (20�–50�). The relationship is robust across experi-

ments and seasons for both the NH and SH, with a

correlation of 0.50, significant at the 99% confidence level.

For the NH only, the correlation coefficient increases to

0.65; similarly, the relationship improves to 0.60 based on

DJF, JJA and ANN only. Similar results are obtained with

other subtropical/midlatitude boundaries. This relationship

indicates that a more positive FNET on the equatorward

flank of the jet is associated with jet expansion

(ANTHRO), or alternatively, that a less positive FNET on

the equatorward flank of the jet is associated with jet

contraction (NAT). One possible explanation is that

expansion of the tropics is a mechanism to increase the heat

transported to midlatitudes. Note, however, that dFNET/dY

explains only 25% of the variance of D/jet, which suggests

other mechanisms–possibly related to stratospheric DT-

may be important.

Similar to what occurs with greenhouse gas (GHG) (e.g.

Kushner et al. 2001; e.g. Raisanen 2003) and IPCC AR4

(e.g. Lorenz and DeWeaver 2007) experiments, both UHT

jets rise and strengthen (not shown). This is consistent with

maximum warming in the tropical upper troposphere,

leading to a higher tropopause and an increase in meridional

Table 3 Poleward displacement of the NH and SH subtropical jet stream (STJ), subtropical mean meridional mass circulation minima (MMC),

and subtropical precipitation minus evaporation (P - E) transition for both SOM (top) and DOM (bottom) ocean boundary conditions

DJF JJA ANN

STJ MMC P - E STJ MMC P - E STJ MMC P - E

SOM

NH

ANTHRO -0.10 0.40** 0.32** 0.20 0.24** 0.27 0.17 0.02 0.27*

NAT -0.18 -0.20* -0.26** -0.22 -0.19* -0.23 -0.16 -0.02 -0.22

UHT 0.28 0.05 0.29 -0.48 -0.43* 0.70 -0.11 -0.23 0.33

SH

ANTHRO -0.12 -0.46** -0.07 0.27 0.13 0.11 -0.02 0.36 -0.02

NAT -0.14 0.39** -0.01 -0.15 -0.13 -0.06 -0.03 -0.06 -0.01

UHT -0.42 -0.48 0.05 -0.46 -0.11 0.37 -0.37 -0.41 0.21*

DOM

NH

ANTHRO -0.12 0.09 0.13 0.28 0.29** -0.49 0.20 -0.30 -0.11

NAT 0.11 0.11 0.06 -0.18 -0.12 -0.26 -0.11 0.20 -0.11

SH

ANTHRO 0.19 -0.09 -0.03 0.05 0.17* -0.05 0.11 0.72** 0.04

NAT 0.08 0.08 0.04 -0.17 -0.06 0.13 0.04 -0.32 0.05

The P - E transition is defined as the latitude where P - E changes from negative to positive on the poleward sides of the the subtropical

minima. Significance is denoted by bold (C90%); *(C95%) and **(C99%). Units are degrees latitude

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

temperature gradient. By the thermal wind equation, this

implies an increase in westerlies. Global warming experi-

ments, however, also feature poleward displacement of the

subtropical jets (Seidel et al. 2008), which is opposite the

contraction that occurs with UHT (Table 3). Because UHT

and GHG experiments feature similar tropospheric warm-

ing patterns, we suspect the dissimilar jet response is due to

the different stratospheric temperature response. We are

currently exploring this GHG–UHT difference, and others,

in more detail.

3.3.2 Mid/high-latitude response

Aerosol also have a significant affect on zonal winds

outside the tropics. Figure 5 shows the ANTHRO DOM

run exhibits a significant increase in wind near 60�N in

DJF throughout the troposphere and lower stratosphere,

with a decrease near 30�N; an opposite pattern is forced

by NAT. Moreover, Fig. 8 shows that NAT and

ANTHRO, particularly for the DOM experiments, exhibit

opposite DJF sea-level pressure changes (DPSL) at high

NH latitudes. For NAT DOM, DPSL is positive poleward

of 50�N, peaking at 1.7 hPa near 90�N. For ANTHRO

DOM, however, DPSL is negative poleward of 50�N,

peaking at -2 hPa. A similar, but even larger, response

occurs during MAM, with DPSL peaking at 1.9 hPa for

NAT DOM and -3 hPa for ANTHRO DOM (not

shown). These changes are reminiscent of the positive

and negative AO (Thompson and Wallace 2000),

respectively, and are associated with DT and the

propagation of wave activity (see next section). The

ANTHRO changes are similar to those reported by

Chung and Ramanathan (2003), who found that Indian

aerosols drove zonal wind changes consistent with the

positive AO.

Fig. 6 Zonal JJA gross dry static stability change (blue;

DhTP � Dhsfc), tropopause pressure change (red; DPTP), climatolog-

ical baroclinicity at 300 hPa (green solid; dT/dY) and the corre-

sponding response (green dashed; DdT=dY) for ANTHRO SOM (top)

and NAT SOM (bottom). Northern hemisphere baroclinicity, and the

response, have been multiplied by -1 so that positive values represent

colder air poleward

Fig. 7 Scatterplot of the poleward jet displacement (D/jet; degrees

latitude) versus the meridional net aerosol forcing gradient (dFNET/

dY) averaged over 20–50� for each aerosol experiment, hemisphere

and season (DJF, JJA, SON, MAM and ANN). NH dFNET/dY values

have been multiplied by -1 so that positive values represent more

positive FNET at lower latitudes. Also shown is the correlation

coefficient (r) and the corresponding best fit line and slope (m). Net

aerosol forcing is estimated as FATM ? FSFC, with surface forcing

over ocean grid points set to zero for the DOM experiments

Fig. 8 DJF sea level pressure change. Units are hPa. Symbolsrepresent significance as in Fig. 2

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

In addition to the troposphere/stratosphere wind

changes, aerosols affect mid/high-latitude winds near the

surface, a measure of the storm tracks. Figure 9 shows

the zonal wind at the model’s bottom level (UBOT) and

the corresponding change, for the SOM experiments

during DJF and JJA. During DJF in the NH, ANTHRO

shifts the storm track poleward, while NAT shifts it

southward, consistent with the above analysis. DOM

results (not shown) are very similar. The corresponding

SOM changes during austral winter (JJA) are less dis-

tinct, but show anthropogenic aerosols yield a

strengthening and slight poleward shift the SH winter-

time storm track, with non-significant changes for nat-

ural aerosols. The ANTHRO results are similar to GHG

experiments of the twenty-first century, which exhibit

intensification and a poleward shift of the storm tracks

in both hemispheres during DJF and in the SH during

JJA (Yin 2005).

4 Understanding the high latitude response

Changes in the AO associated with changed climates have

previously been explained by changes in wave propagation

and/or refractive index of the atmosphere (e.g., Rind et al.

2005). Under quasi-geostrophic scaling, the northward (Fy)

and vertical (Fz) Eliassen-Palm (EP) flux (F), which are

proportional to the meridional and vertical zonal average

wave energy flux (Holton 1992), is

Fy ¼ �qu0v0; Fz ¼qfRdv0T 0

N2Hð1Þ

where u0v0 is the meridional eddy momentum flux, v0T 0 is

the eddy northward heat flux, N is the Brunt-Vaisala fre-

quency, Rd is the dry air gas constant, H is the scale height,

f is the Coriolis parameter and q is the air density.

Figure 10 shows the DJF change in EP flux and flux

divergence, divided by the standard density. The changes

are largest for the DOM (fixed-SST) signals, with consis-

tently opposite changes for NAT and ANTHRO forcings.

NAT shows upward and poleward wave energy propaga-

tion, peaking between 40 and 60�N. This is consistent with

a decrease in wave refraction, defined as the ratio of the

meridional and vertical EP flux [the meridional turning

relative to the upward energy flux, as in Rind et al. (2005)].

The average NH percent change in wave refraction is

-6.5% for NAT DOM and -3.7% for NAT SOM (18 and

10.2%, respectively, for ANTHRO). NAT forcing also

yields anomalous F convergence above 200 hPa and

poleward of 40�N. This suggests an increase in westward

momentum deposition (i.e., wave drag), inducing a torque

that decelerates the winds. ANTHRO changes are again

consistently opposite to these.

A spectral analysis indicates that DFz arises mainly from

zonal wavenumbers 1–4 (Fig. 11). For both DOM runs,

DFz is mainly composed of zonal wavenumber 2, while

for SOM, the response is dominated by wavenumber 3 for

NAT forcing and shows no dominant component for

ANTHRO forcing. By the Charney-Drazin theory (e.g.,

Andrews et al. 1987), the horizontal length scale of plan-

etary waves controls the extent to which they can propagate

into the stratosphere. This suggests that wave energy was

better able to penetrate into the stratosphere in the DOM

experiments due to the longer wavelength of the response,

explaining the stronger AO-like changes near the surface in

those runs.

The reason for the prominence of wavenumber 2 in the

DOM response is likely because this wavenumber is also

prominent in the net aerosol forcing at mid-latitudes

(30–60�N), where in DOM, net forcing over oceans

includes only the atmospheric heating part. The DOM net

forcing is near zero over the Pacific and Atlantic Oceans,

but has negative maxima at two approximately opposite

longitudes: in the India/SE Asia and North America

regions. In the SOM experiments, however, the net aerosol

forcings are negative over both oceans and land. This is

especially true for ANTHRO because of the more com-

plicated geographical pattern of the anthropogenic aerosol

amount and optical properties.

Fig. 9 As in Fig. 4, except for zonal wind at the model’s bottom

level (UBOT)

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

The increased wave drag in NAT is associated with

increased poleward transport of air in the middle strato-

sphere and subsidence in the lower stratosphere. This

hydrostatically leads to increased mass and PSL at high

latitudes (Fig. 8). The increased subsidence also leads to

adiabatic compression, increasing temperatures in the

lower stratosphere at high latitudes (Fig. 2). By thermal

wind balance, this reduced meridional temperature gradient

implies a decrease in the strength of the westerlies in the

mid-stratosphere (Fig. 5). This weakened shear suggests

less refraction of planetary wave energy. Since planetary

waves are propagating up from the surface, there is less

refraction by the decreased vertical shear below the area of

reduced zonal wind. Less refraction of planetary wave

energy at the lower boundary of the wind anomaly leads to

wave convergence (Fig. 10) and a deceleration of the zonal

wind in that area. Over time, the wind anomaly propagates

downward from the mid-stratosphere to the troposphere

(Haynes et al. 1991; Shindell et al. 2001; Stenchikov et al.

2002; Song and Robinson 2004). Opposite changes occur

with ANTHRO.

The EP flux changes are consistent with the tropo-

spheric temperature changes, where an increase in

rT suggests an increase in mean zonal energy and

amplitude of planetary waves (Stenchikov et al. 2002).

For DOM, both ANTHRO and NAT forcings warm the

tropical troposphere due to the imposed aerosol heating.

However, NAT cools near 40�N, resulting in a stronger

meridional rT; ANTHRO warms most near 40�N,

resulting in a weaker meridional rT. Similar meridional

DrT exist for the SOM signals.

Figure 12 shows the changes in geopotential height,

temperature and horizontal winds at 500 hPa for both DOM

runs. The aerosol heating induces local circulation changes

in the mid-latitudes, concentrated into four centers near

40�N: the central Pacific Ocean, North America, the east-

ern Atlantic Ocean and northeast China/Japan (these cen-

ters also occur in the SOM experiments). The latter two

teleconnections resemble those Kim et al. (2006) find

during MAM. With NAT forcing these centers are asso-

ciated with lower geopotential heights, counterclockwise

circulation anomalies and colder temperatures. The latter

are associated with both cold-air advection and adiabatic

cooling of ascending air (not shown). Midlatitude cooling

yields a greater rT, which is associated with increased

zonal mean energy and amplitudes of tropospheric plane-

tary wave, and increased wave flux activity into the lower

stratosphere. These results are similar to those of

Stenchikov et al. (2002), where the tropospheric cooling

effect of volcanic aerosols decreases rT between 30 and

60�N, resulting in decreased vertical wave activity flux and

circulation changes consistent with the positive AO.

Fig. 10 DJF EP flux (m2 s-2)

and EP flux divergence (10-6 m

s-2) change, divided by the

standard density. Contour

interval is {-6, -4, -2,

-1, 0, 1, 2, 4, 6}

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

Our attribution of the AO-like changes to the spatial

heterogeneity of the aerosol forcing—and resulting low/

mid latitude DrT—is also supported by the changes in the

UHT experiment, where changes in low/mid-latitude

rT and the AO were both negligible. As mentioned above,

however, the vertical profile of UHT warming in the tro-

posphere is similar to GHG experiments, which generally

exhibit positive AO changes (Shindell et al. 2001; Gillett

et al. 2002; Osborn 2004; Rauthe et al. 2004), as do IPCC

models with both GHGs and sulfate aerosols (Miller et al.

2006). Such an AO trend is generally explained by an

increase in stratospheric rT, which produces stronger mid-

latitude zonal winds that tend to inhibit propagation of

planetary waves into the polar stratosphere (Stenchikov

et al. 2002). The lack of a UHT AO signal, therefore,

appears to be related to the dissimilar zonal wind response,

with negligible increases in mid-latitude UHT U during

DJF.

5 Conclusions

We used the Community Atmosphere Model (CAM3) to

investigate the equilibrium climate response to anthropo-

genic and natural aerosols, focusing on atmospheric cir-

culation changes. To avoid the uncertainties and known

underestimation of directly simulated absorbing aerosol

forcing (Sato et al. 2003; Ramanathan and Carmichael

2008; Koch et al. 2009), we directly incorporated the

aerosol direct effect (i.e., absorption and reflection of

solar radiation) estimated from the model-assisted data

analysis of Chung et al. (2005). Their estimated top-

of-atmosphere, direct, clear-sky forcing (3.4 W m-2) is

nearly three times the average (1.24 W m-2) reported

from GCMs by Forster et al. (2007). To help determine

the importance of oceans in communicating forcing

responses back to the atmosphere we repeated each

forcing experiment with both climatological SSTs and a

slab ocean model. Results show that aerosols affect the

circulation at all latitudes, with natural (mostly reflecting)

and anthropogenic (mostly absorbing) aerosols driving

changes to the circulation that are qualitatively similar but

of opposite sign. Despite the geographically uneven

forcing, the planetary-scale responses appear to be

attributable to the zonal-mean (or in some cases low-

wavenumber) aspects of the forcing.

An important caveat to our study is that anthropogenic

aerosols are a mixture of at least two types (sulfate and

black carbon) that have very different optical properties,

and at present seem to occur in a proportion so as to exert a

relatively small net top-of-atmosphere radiative effect. We

anticipate based on our results that anthropogenic sulfate

would, by itself, have planetary-scale impacts similar to

those of natural aerosols (both having high single-scatter

albedo) despite their somewhat different regional patterns.

Given the above, and the finding here of nearly opposite

effects of anthropogenic and natural aerosols, we infer that

black carbon alone would have significantly stronger

impacts than found in our anthropogenic aerosol experi-

ments. Since the relative amounts of sulfate and black

carbon have changed significantly over time, decadal

variations in aerosol forcing and response may well have

Fig. 11 DJF change in the

vertical component of the EP

flux (10-3 m2 s-2) at 600 hPa

divided by the standard density

(black line). Also included is the

contribution to DFz from each

of the four leading zonal

wavenumbers (colored lines)

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

exceeded those for either the present-day anthropogenic or

natural aerosol. Further studies would be needed to test

these assertions.

Due to the predominance of northern-hemisphere sour-

ces for both aerosol types considered, anthropogenic

aerosols warmed the troposphere (and natural aerosols

cooled it) more in the northern hemisphere than in the

southern, with changes of order 0.1–0.3 K in the lower

troposphere. Anthropogenic aerosols consequently shifted

the ITCZ northward while natural aerosols shifted it

southward. The northward shift is associated with a

weakening of the DJF mean meridional mass circulation

and strengthening of the JJA one, with opposite changes for

the southward shift; all are consistent with the radiatively

forced changes to inter-hemispheric temperature gradients.

This behavior is consistent with other aerosol studies

focusing on the direct effects of BC aerosols (Roberts and

Jones 2004; Wang 2004, 2007; Chung and Seinfeld 2005;

Yoshimori and Broccoli 2008), and the direct (Yoshimori

and Broccoli 2008) and indirect (Rotstayn et al. 2000;

Williams et al. 2001) effects of sulfate aerosols. Changes

in Hadley cell strength were smaller in the fixed-SST

experiments because inter-hemispheric temperature gradi-

ents were not able to change as much. These results support

previous findings that aerosols affect the variability of

precipitation at low latitudes, for example in the Amazon

(Cox et al. 2008) and the Sahel (Rotstayn and Lohmann

2002). Aerosol forcing is also associated with meridional

shifts of the subtropical jets. In the slab-ocean experiments,

anthropogenic aerosols move the subtropical jets poleward

by 0.2�–0.3� each, leading to expansion of the tropics.

Natural aerosols produce the opposite effect.

Global emissions of black carbon have generally

increased over the latter half of the twentieth century,

although they remain quite uncertain and have probably

fallen somewhat since 1990 (Novakov et al. 2003; Ito and

Fig. 12 DJF geopotential height (left panel) and temperature (rightpanel) change at 500 hPa. Horizontal wind change (m s-1) is included

with DT: Symbols—which are plotted every other grid point to

improve clarity—represent significance at the 90% (diamond), 95%

(cross) and 99% (dot) confidence level. Units are m and K,

respectively

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

Penner 2005; Bond et al. 2007). Global emissions of sul-

fate aerosols, however, have been declining since the 1970s

(van Aardenne et al. 2001; Smith et al. 2004). Our results

indicate that both of these trends should have contributed to

poleward migration of the subtropical jet in the NH, and

possibly in the SH, hence contributing to the observed

widening of the tropics from the 1970s through 1990 or so.

In fact such widening has been observed (or inferred from

stratospheric cooling and tropospheric warming trends),

and is larger than predicted by models forced with GHGs

and other forcings (Fu et al. 2006; Seidel et al. 2008; Jo-

hanson and Fu 2009). Although some of these models

include aerosol forcing, aerosol absorption is likely

underestimated (Sato et al. 2003; Koch et al. 2009). The

observed widening of 2.0�–4.8� over 25 years, however, is

much larger than reported here for either aerosol forcing

(*0.5�) and does not appear to have stopped in the last

decade or two. Nonetheless, aerosols may have contributed

non-negligibly to this widening and, as discussed above,

impacts from past changes in anthropogenic aerosol com-

position could exceed those simulated here for the current

composition.

Arctic oscillation-like changes result from altered tro-

pospheric temperature gradients, which affect the vertical

propagation of wave activity. We argue that this is because

anthropogenic aerosols decrease temperature gradients

between low and mid-latitudes, decreasing the vertically

propagating wave activity and increasing equatorward

refraction, with opposite impacts from natural aerosols.

The increased refraction causes acceleration of the strato-

spheric zonal winds, which eventually propagates back

down through the troposphere (Haynes et al. 1991; Shin-

dell et al. 2001; Stenchikov et al. 2002; Song and Robin-

son 2004) where it manifests itself at the surface as sea-

level pressure and temperature anomalies.

The result is zonal winds near 60�N increasing by *1 m

s-1, temperatures in the high-latitude stratosphere decreas-

ing by (*1 K) and high-latitude sea-level pressure

decreasing by *2 hPa, with anthropogenic aerosol forcing.

Similar impacts occur in the simulation of Chung and Ra-

manathan (2003) for absorbing aerosols over India only. We

found that changes were significant only with fixed SSTs,

apparently because longer wavelength planetary waves—

which are better able to penetrate into the stratosphere—are

preferentially excited by the imposed net aerosol forcing in

this case due to the land–ocean distribution in the northern

hemisphere. Regionally restricted forcings could excite a

similar response even with interactive oceans.

Because the high-latitude AO impacts are strong only

with fixed SSTs, they do not appear to be robust to varia-

tions in ocean behavior, and fixed-SST results are unlikely

to represent very well the impacts of trends in aerosols

where the ocean has plenty of time to respond to flux

changes at the surface. Moreover, the observed changes are

significantly larger than those reported here even with fixed

SST: from 1965 to 1995, mean sea-level pressure north of

45�N dropped by 2.5 hPa relative to that from 45�N to the

equator (Gillett 2005), compared with a peak response here

of 0.4 hPa. Similarly, zonal wind increased by 7 m s-1 at

60�N and 50 hPa (Scaife et al. 2005), compared to roughly

1 m s-1 here. Thus, we find wind and pressure changes that

occur in roughly the same ratio as those of recent hard-to-

explain trends, but at much smaller magnitudes.

Nonetheless, the decadal variability in aerosol forcing

(e.g., SE Asian haze Ramanathan et al. 2001b; e.g., Chung

and Ramanathan 2003)—as opposed to, say, the more

monotonically changing forcing by greenhouse gases—

makes it an interesting possibility for explaining variations in

the AO, which also have a strong decadal nature (Feldstein

2002). Given the cancellation found here between absorbing

and scattering aerosol impacts, it is possible that decadal

changes in the ratio of black carbon to sulfate could have

exerted large effects. It is also possible that shifts of emis-

sions from one region to another (Streets et al. 2009) may

have affected the AO by influencingrT and the wavelength

of perturbations to the midlatitude flow. It would appear

worthwhile to include more realistic aerosol forcing changes

in climate models, or at least to consider more seriously the

possible impacts of unknown variations in the distribution

and type of aerosols as an additional source of forcing

uncertainty in model experiments.

Acknowledgments This study was funded by the NSF (CAREER

ATM-0453639), Yale University, and by NSF ARC-0714088 and

NASA NNX07AR23G, UC Irvine. We thank Chul Eddy Chung for

kindly providing his aerosol data set and for other useful suggestions,

Charlie Zender for providing computing resources at UC Irvine, and

Dorothy Koch and V. Ramanathan for useful suggestions. Part of this

manuscript contributed to RJA’s PhD thesis.

Open Access This article is distributed under the terms of the

Creative Commons Attribution Noncommercial License which per-

mits any noncommercial use, distribution, and reproduction in any

medium, provided the original author(s) and source are credited.

References

Ackerman AS, Toon OB, Stevens DE, Heymsfield AJ, Ramanathan

V, Welton EJ (2000) Reduction of tropical cloudiness by soot.

Science 288:1042–1047

Albrecht BA (1989) Aerosols, cloud microphysics, and fractional

cloudiness. Science 245:1227–1230

Allen RJ, Sherwood SC (2010) The aerosol-cloud semi-direct effect

and land-sea temperature contrast in a GCM. Geophys Res Lett.

37:L07702. doi:10.1029/2010GL042759

Anderson TL, Wu Y, Chu DA, Schmid B, Redemann J, Dubovik O

(2005) Testing the MODIS satellite retrieval of aerosol fine-

mode fraction. J Geophys Res 110(D18204). doi:10/1029/2005

JD005978

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

Andrews DG, Holton JR, Leovy CB (1987) Middle atmosphere

dynamics. Elsevier, New York, pp 489

Bellouin N, Boucher O, Haywood J, Reddy MS (2005) Global

estimates of aerosol direct radiative forcing from satellite

measurements. Nature 438:1138–1141

Bollasina M, Nigam S, Lau K-M (2008) Absorbing aerosols and

summer monsoon evolution over South Asia: an observational

portrayal. J Clim 21:3221–3239

Bond TC, Bhardwaj E, Dong R, Jogani R, Jung S, Roden C, Streets

DG, Trautmann NM (2007) Historical emissions of black and

organic carbon aerosol from energy related combustion,

1850–2000. Global Biogeochem Cycles 21(GB2018). doi:

10.1029/2006GB002840

Broccoli AJ, Dahl KA, Stouffer RJ (2006) Response of the ITCZ to

Northern Hemisphere cooling. Geophys Res Lett 33:L01702.

doi:10.1029/2005GL024546

Chin M et al (2002) Tropospheric aerosol optical thickness from the

GOCART model and comparisons with satellite and sunpho-

tometer measurements. J Atmos Sci 59:461–483

Chung C, Ramanathan V (2006) Weakening of the North Indian SST

gradients and the monsoon rainfall in India and the Sahel. J Clim

19:2036–2045

Chung CE (2006) Steady vs. fluctuating aerosol radiative forcing in a

climate model. J Korean Meteor Soc 42(6):411–417

Chung CE, Ramanathan V (2003) South Asian haze forcing: remote

impacts with implications to ENSO and AO. J Clim

16:1791–1806

Chung CE, Ramanathan V, Kiehl JT (2002) Effects of the south Asian

absorbing haze on the northeast monsoon and surface-air heat

exchange. J Clim 15:2462–2476

Chung CE, Ramanathan V, Kim D, Podgorny IA (2005) Global

anthropogenic aerosol direct forcing derived from satellite and

ground-based observations. J Geophys Res 110(D24207). doi:

10.1029/2005JD006356

Chung SH, Seinfeld JH (2005) Climate response of direct radiative

forcing of anthropogenic black carbon. J Geophys Res

110(D11):D11102. doi:10.1029/2004JD005441

Collins WD et al (2004) Description of the NCAR Community

Atmosphere Model (CAM 3.0). Tech. Rep. NCAR/TN-

464?STR, National Center for Atmospheric Research, Boulder,

CO, 226 pp

Cox PM et al (2008) Increasing risk of Amazonian drought due to

decreasing aerosol pollution. Nature 453. doi:10.1038/nature

06960

Evan AT, Vimont DJ, Heidinger AK, Kossin JP, Bennartz R (2009)

The role of aerosols in the evolution of tropical North Atlantic

Ocean temperature anomalies. Science 324:778–781

Feldstein SB (2002) The recent trend and variance increase of the

annular mode. J Clim 15:88–94

Forster P et al (2007) Changes in atmospheric constituents and in

radiative forcing. In: Climate change 2007: the physical science

basis. Contribution of working group I to the fourth assessment

report of the intergovernmental panel on climate change.

Cambridge University Press, Cambridge

Frierson DMW, Lu J, Chen G (2007) Width of the Hadley cell in

simple and comprehensive general circulation models. Geophys

Res Lett 34:L18804. doi:10.1029/2007GL031115

Fu Q, Johanson CM, Wallace JM, Reichler T (2006) Enhanced mid-

latitude tropospheric warming in satellite measurements. Science

312:1179

Gillett NP (2005) Northern Hemisphere circulation. Nature

437:496–496

Gillett NP, Allen MR, McDonald RE, Senior CA, Shindell DT,

Schmidt GA (2002) How linear is the Arctic Oscillation response

to greenhouse gases? J Geophys Res 107:D3. doi:10.1029/2001

JD000589

Haigh JD, Blackburn M, Day R (2005) The response of tropospheric

circulation to perturbations in lower-stratospheric temperature.

J Clim 18:3672–3685

Haynes PH, Marks CJ, McIntyre ME, Shepherd TG, Shine KP (1991)

On the ‘‘downward control’’ of extratropical diabatic circulations

by eddy-induced mean zonal forces. J Atmos Sci 48:651–678

Held IM (2000) The general circulation of the atmosphere, paper

presented at 2000 Woods Hole Oceanographic Institute Geo-

physical Fluid Dynamics Program, Woods Hole Oceanographic

Institute, Woods Hole, MA. http://www.whoi.edu/page.do?

pid=13076

Holton JR (1992) An introduction to dynamic meteorology.

Academic Press, New York, 511 pp

Hu Y, Fu Q (2007) Observed poleward expansion of the Hadley

circulation since 1979. Atmos Chem Phys. 7:5229–5236

Ito A, Penner JE (2005) Historical emissions of carbonaceous aerosols

from biomass and fossil fuel burning for the period 1870–2000.

Global Biogeochem Cycles 19(GB2028). doi:10.1029/2004

GB002374

Johanson CM, Fu Q (2009) Hadley cell widening: model simulations

versus observations. J Clim 22:713–2725

Kim MK, Lau WKM, Chin M, Kim KM, Sud YC, Walker GK (2006)

Atmospheric teleconnection over Eurasia induced by aerosol

radiative forcing during boreal spring. J Clim 19:4700–4718

Koch D et al (2009) Evaluation of black carbon estimations in global

aerosol models. Atmos Chem Phys 9(22):9001–9026

Kushner PJ, Held IM, Delworth TL (2001) Southern Hemisphere

atmospheric circulation response to global warming. J Clim

14:2238–2249

Levy RC et al (2003) Evaluation of the moderate-resolution imaging

spectroradiometer (MODIS) retrievals of dust aerosol over the

ocean during PRIDE. J Geophys Res 108(D19). doi:

10.1029/2002JD002460

Lorenz DJ, DeWeaver ET (2007) Tropopause height and zonal wind

response to global warming in the IPCC scenario integrations.

J Geophys Res 112. doi:10.1029/2006JD008087

Lu J, Deser C, Reichler T (2009) Cause of the widening of the tropical

belt since 1958. Geophys Res Lett 36:L03803. doi:

10.1029/2008GL036076

Lu J, Vecchi GA, Reichler T (2007) Expansion of the Hadley cell

under global warming. Geophys Res Lett 34:L06805. doi:

10.1029/2006GL028443

Mantsis DF, Clement AC (2009) Simulated variability in the mean

atmospheric meridional circulation over the 20th century.

Geophys Res Lett 36:L06704. doi:10.1029/2008GL036741

Meehl GA, Arblaster JM, Collins WD (2008) Effects of black carbon

aerosols on the Indian monsoon. J Clim 21:2869–2882

Menon S, Hansen J, Nazarenko L, Luo Y, (2002) Climate effects of

black carbon aerosols in China and India. Science 297:2250–2253

Miller RL, Schmidt GA, Shindell DT (2006) Forced annular

variations in the 20th century Intergovernmental Panel on

Climate Change Fourth Assessment Report models. J Geophys

Res 111:D18101. doi:10.1029/2005JD006323

Mitchell JFB, Johns TC, Gregory JM, Tett SFB (1995) Climate

response to increasing levels of greenhouse gases and sulphate

aerosols. Nature 370:501–504

Novakov T, Ramanathan V, Hansen JE, Kirchstetter TW, Sato M,

Sinton JE, Sathaye JA (2003) Large historical changes of fossil-

fuel black carbon aerosols. Geophys Res Lett 30(6). doi:

10.1029/2002GL016345

Osborn TJ (2004) Simulating the winter North Atlantic Oscillation:

the roles of internal variability and greenhouse gas forcing. Clim

Dyn 22:605–623

Podgorny IA, Conant WC, Ramanathan V, Satheesh SK (2000)

Aerosol modulation of atmospheric and solar heating over the

tropical Indian Ocean. Tellus B 52:947–958

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123

Raisanen J (2003) CO2-induced changes in the atmospheric angular

momentum in CMIP2 experiments. J Clim 16:132–143

Ramanathan V, Carmichael G (2008) Global and regional climate

changes due to black carbon. Nat Geosci 1:221–227

Ramanathan V, Crutzen PJ, Kiehl JT, Rosenfeld D (2001a) Aerosols,

climate and the hydrological cycle. Science 294:2119–2124

Ramanathan V, Crutzen PJ, Lelieveld J, Mitra AP et al. (2001b)

Indian Ocean Experiment: an integrated analysis of the climate

forcing and effects of the great Indo-Asian haze. J Geophys Res

106(D22):28371–28398

Ramanathan V, Ramana MV, Roberts G, Kim D, Corrigan C, Chung

C, Winker D (2007) Warming trends in Asia amplified by brown

cloud solar absorption. Nature 448:575–578

Ramanathan V et al (2005) Atmospheric brown clouds: impacts on

South Asian climate and hydrological cycle. Proc Natl Acad Sci

102:5326–5333

Randles CA, Ramaswamy V (2008) Absorbing aerosols over Asia: a

geophysical fluid dynamics laboratory general circulation model

sensitivity study of model response to aerosol optical depth and

aerosol absorption. J Geophys Res 113:D21203. doi:

10.1029/2008JD010140

Rauthe M, Hense A, Paeth H (2004) A model intercomparison study

of climate change signals in extratropical circulation. Int J

Climatol 24:643–662

Reichler T, Dameris M, Sausen R (2003) Determining the tropopause

height from gridded data. Geophys Res Lett 30(20):2042. doi:

10.1029/2003GL018240

Rind D, Perlwitz J, Lonergan P (2005) AO/NAO response to climate

change: 1. Respective influences of stratospheric and tropo-

spheric climate changes. J Geophys Res 110(D12107). doi:

10.1029/2004JD005103

Roberts DL, Jones A (2004) Climate sensitivity to black carbon

aerosol from fossil fuel combustion. J Geophys Res 109:D16202.

doi:10.1029/2004JD004676

Robock A (2000) Volcanic eruptions and climate. Rev Geophys

38:191–219

Roeckner EL, Bengtsson L, Feichter J, Lelieveld J, Rodhe H (1999)

Transient climate change simulations with a coupled atmo-

sphere–ocean GCM including the tropospheric sulfur cycle.

J Clim 12:3004–3031

Rotstayn LD, Lohmann U (2002) Tropical rainfall trends and the

indirect aerosol effect. J Clim 15:2103–2116

Rotstayn LD, Ryan BF, Penner JE (2000) Precipitation changes in a

GCM resulting from the indirect effects of anthropogenic

aerosols. Geophys Res Lett 27(19):3045–3048

Sato M et al (2003) Global atmospheric black carbon inferred from

AERONET. Proc Natl Acad Sci 100(11):6319–6324

Scaife AJ, Knight JR, Vallis GK, Folland CK (2005) A stratospheric

influence on the winter NAO and North Atlantic surface climate.

Geophys Res Lett 32:L18715. doi:10.1029/2005GL023226

Seidel DJ, Fu Q, Randel WJ, Reichler TJ (2008) Widening of the

tropical belt in a changing climate. Nat Geosci 1. doi:

10.1038/ngeo.2007.38

Shindell DT, Faluvegi G (2009) Climate response to regional

radiative forcing during the twentieth century. Nat Geosci. 2.

doi:10.1038/NGEO473

Shindell DT, Schmidt GA, Mann ME, Faluvegi G (2004) Dynamic

winter climate response to large tropical volcanic eruptions since

1600. J Geophys Res 109:D05104. doi:10.1029/2003JD004151

Shindell DT, Schmidt GA, Miller RL, Rind D (2001) Northern

Hemisphere winter climate response to greenhouse gas, ozone,

solar, and volcanic forcing. J Geophys Res 106(D7):7193–7210

Smith S, Andres R, Conception E, Lurz J (2004) Historical sulfur

dioxide emissions 1850–2000: methods and results. Tech. Rep.

PNNL-14537, Joint Global Change Research Institute

Song Y, Robinson WA (2004) Dynamical mechanisms for strato-

spheric influences on the troposphere. J Atmos Sci

61:1711–1725

Stenchikov G, Robock A, Ramaswamy V, Schwarzkopf MD,

Hamilton K, Ramachandran S (2002) Arctic Oscillation response

to the 1991 Mount Pinatubo eruption: effects of volcanic

aerosols and ozone depletion. J Geophys Res 107(D24):4803.

doi:10.1029/2002JD002090

Streets DG et al (2009) Anthropogenic and natural contributions to

regional trends in aerosol optical depth, 1980–2006. J Geophys

Res 114(D00D18). doi:10.1029/2008JD011624

Thompson DWJ, Wallace JM (2000) Annular modes in the

extratropical circulation. Part I: Month to month variability.

J Clim 13:1000–1016

Twomey S (1977) The influence of pollution on the shortwave albedo

of clouds. J Atmos Sci 34:1149–1152

van Aardenne JA, Dentener FJ, Olivier JGJ, Klein Goldewijk CGM,

Lelieveld J (2001) A 1 x 1 degree resolution dataset of historical

anthropogenic trace gas emissions for the period 1890–1990.

Global Biogeochem Cycles 15(4):909–928

Wang C (2004) A modeling study on the climate impacts of black

carbon aerosols. J Geophys Res 109:D03106. doi:

10.1029/2003JD004084

Wang C (2007) Impact of direct radiative forcing of black carbon

aerosols on tropical convective precipitation. Geophys Res Lett

34:L05709. doi:10.1029/2006GL028416

Williams GP (2006) Circulation sensitivity to tropopause height.

J Atmos Sci 63:1954–1961

Williams KD, Jones A, Roberts DL, Senior CA, Woodage MJ (2001)

The response of the climate system to the indirect effects of

anthropogenic sulfate aerosol. Clim Dyn 17:845–856

Yin JH (2005) A consistent poleward shift of the strom tracks in

simulations of 21st century climate. Geophys Res Lett

32:L18701. doi:10.1029/2005GL023684

Yoshimori M, Broccoli AJ (2008) Equilibrium response of an

atmosphere-mixed layer ocean model to different radiative

forcing agents: global and zonal mean response. J Clim

21:4399–4423

Yu H et al (2006) A review of measurement-based assessments of the

aerosol direct radiative effect and forcing. Atmos Chem Phys6:613–666

Zhang R, Delworth TL (2005) Simulated tropical response to a

substantial weakening of the Atlantic thermohaline circulation.

J Clim 18:1853–1860

R. J. Allen, S. C. Sherwood: The impact of natural versus anthropogenic aerosols

123